Numerical age and geologic time

Dating Sedimentary Rocks? 

The mind grows giddy gazing so far back into the abyss of time. John Playfair (1747–1819),  British geologist who popularized the works of Hutton.

We have seen that isotopic dating can be used to date the time when igneous rocks formed and when metamorphic rocks metamorphosed, but not when sedimentary rocks were deposited. So how do we determine the numerical age of a sedimentary rock? We must answer this question if we want to add numerical ages to the geologic column. Geologists obtain dates for sedimentary rocks by studying cross-cutting relationships between sedimentary rocks and datable igneous or metamorphic rocks. For example, if we find a sequence of sedimentary strata deposited unconformably on a datable granite, the strata must be younger than the granite  (figure above). If a datable basalt dike cuts the strata, the strata must be older than the dike. And if a datable volcanic ash buried the strata, then the strata must be older than the ash.

The Geologic Time Scale 

Geologists have searched the world for localities where they can recognize cross-cutting relations between datable igneous  

rocks and sedimentary rocks or for layers of datable volcanic rocks inter-bedded with sedimentary rocks. By isotopically dating the igneous rocks, they have been able to provide numerical ages for the boundaries between all geologic periods. For example, work from around the world shows that the Cretaceous Period began about 145 million years ago and ended 65 million years ago. So the Cretaceous sandstone bed in first figure was deposited during the middle part of the Cretaceous, not at the beginning or end. 

The discovery of new data may cause the numbers defining the boundaries of periods to change, which is why the term numerical age is preferred to absolute age. In fact, around 1995, new dates on rhyolite ash layers above and below the Cambrian-Precambrian boundary showed that this boundary occurred at 542 million years ago, in contrast to previous, less definitive studies that had placed the boundary at 570 million years ago. Figure above shows the currently favoured numerical ages of periods and eras in the geologic column as of 2009. This dated column is commonly called the geologic time scale. 

What Is the Age of the Earth? 

During the 18th and 19th centuries, before the discovery of isotopic dating, scientists came up with a great variety of clever solutions to the question, “How old is the Earth?”—all of which have since been proven wrong. Lord William Kelvin, a 19th century physicist renowned for his discoveries in thermodynamics, made the most influential scientific estimate of the Earth’s age of his time. Kelvin calculated how long it would take for the Earth to cool down from a temperature as hot as the Sun’s, and concluded that this planet is about 20 million years old. Kelvin’s estimate contrasted with those being promoted by followers of Hutton, Lyell, and Darwin, who argued that if the concepts of uniformitarianism and evolution were correct, the Earth must be much older. They argued that physical processes that shape the Earth and form its rocks, as well as the process of natural selection that yields the diversity of species, all take a very long time. Geologists and physicists continued to debate the age issue for many years. The route to a solution didn't appear until 1896, when Henri Becquerel announced the discovery of radioactivity. Geologists immediately realized that the Earth’s interior was producing heat from the decay of radioactive material. This realization uncovered one of the flaws in Kelvin’s argument: Kelvin had assumed that no new heat was produced after the Earth first formed. Because radioactivity constantly generates new heat in the Earth, the planet has cooled down much more slowly than Kelvin had calculated and could be much older. The discovery of radioactivity not only invalidated Kelvin’s estimate of the Earth’s age, it also led to the development of isotopic dating. Since the 1950s, geologists have scoured the planet to identify its oldest rocks. Rocks younger than 3.85 Ga are fairly common. Rock samples from several localities (Wyoming, Canada, Greenland, and China) have yielded dates as old as 4.03 Ga. (Recall that “Ga” means “billion years ago.”) Individual clastic grains of the mineral zircon have yielded dates of up to 4.4 Ga, indicating that rock as old as 4.4 Ga did once exist. Isotopic dating of Moon rocks yields dates of up to 4.50 Ga, and dates on meteorites have yielded ages as old as 4.57 Ga. Geologists consider 4.57-Ga meteorites to be fragments of planetesimals like those from which the Earth first formed. Thus, these dates are close to the age of the Earth’s birth, for models of the Earth’s formation assume that all objects in the Solar System developed at roughly the same time from the same nebula. Why don’t we find rocks with ages between 4.03 and 4.57 Ga in the Earth’s crust? Geologists have come up with several ideas to explain the lack of extremely old rocks. One idea comes from calculations defining how the temperature of our planet has changed over time. These calculations indicate that during the first half-billion years of its existence, the Earth might have been so hot that rocks in the crust remained above the closure temperature for minerals, and isotopic clocks could not start “ticking.” Another idea comes from studies of cratering events on other moons and planets. These studies indicate that the inner planets were bombarded so intensely by meteorites at about 4.0 Ga that almost all crust formed earlier than 4.0 Ga was completely destroyed.

Picturing Geologic Time 

The number 4.57 billion is so staggeringly large that we can’t begin to comprehend it. If you lined up this many pennies in a row, they would make an 87,400-km-long line that would wrap around the Earth’s equator more than twice. Notably, at the scale of our penny chain, human history is only about 100 city blocks long. Another way to grasp the immensity of geologic time is to equate the entire 4.57 billion years to a single calendar year. On this scale, the oldest rocks preserved on Earth date from early February, and the first bacteria appear in the ocean on February 21. The first Shelly invertebrates appear on October 25, and the first amphibians crawl out onto land on November 20. On December 7, the continents coalesce into the super-continent of Pangaea. Birds and the ancestors of mammals  appear about December 15, along with the dinosaurs, and the Age of Dinosaurs ends on December 25. The last week of December represents the last 65 million years of Earth history, including the entire Age of Mammals. The first human-like ancestor appears on December 31 at 3  p.m., and our species, Homo sapiens, shows up an hour before midnight. The last ice age ends a minute before midnight, and all of recorded human history takes place in the last  30 seconds. To put it another way, human history occupies the last 0.000001% of Earth history. The Earth is so old that there has been more than enough time for the rocks and life forms of Earth to have formed and evolved.

How do we determine numerical age of Earth?

Numerical age determination

Geologists since the days of Hutton could determine the relative ages of geologic events, but they had no way to specify numerical ages (called “absolute ages” in older literature). Thus, they could not define a timeline for Earth history or determine the duration of events. This situation changed with the discovery of radioactivity. Simply put, radioactive elements decay at a constant rate that can be measured in the lab and can be specified in years. In the 1950s, geologists developed techniques for using measurements of radioactive elements to calculate the numerical ages of rocks. Geologists originally referred to these techniques as radiometric dating; more recently, this has come to be known as isotopic dating. The overall study of numerical ages is geochronology. Since the 1950s, isotopic dating techniques have steadily improved, and geologists have learned how to make very accurate measurements from very small samples. But the basis of the techniques remains the same, and to explain them, we must first review radioactive decay. 

Radioactive Decay 

All atoms of a given element have the same number of protons in their nucleus we call this number the atomic number. However, not all atoms have the same number of neutrons in their nucleus. Therefore, not all atoms of a given element have the same atomic weight (roughly, the number of protons plus neutrons). Different versions of an element, called isotopes of the element, have the same atomic number but a different atomic weight. For example, all uranium atoms have 92 protons, but the uranium-238 isotope (abbreviated 238U) has an atomic weight of 238 and thus has 146 neutrons, whereas the 235U isotope has an atomic weight of 235 and thus has 143 neutrons. Some isotopes of some elements are stable, meaning that they last essentially forever. Radioactive isotopes are unstable in that eventually, they undergo a change called radioactive decay, which converts them to a different element. Radioactive decay can take place by a variety of reactions that change the atomic number of the nucleus and thus form a different element. In these reactions, the isotope that undergoes decay is the parent isotope, while the decay product is the daughter isotope. For example, rubidium-87 (87Rb) decays to strontium-87 (87Sr), potassium-40 (40K) decays to argon-40 (40Ar), and uranium-238 (238U) decays to lead-206 (206Pb). In some cases, decay takes many steps before yielding a stable daughter. Physicists cannot specify how long an individual radioactive isotope will survive before it decays, but they can measure how long it takes for half of a group of parent isotopes to decay. This time is called the half-life of the isotope. 

Figure above (a-c) can help you visualize the concept of a half-life. Imagine a crystal containing 16 radioactive parent isotopes. (In real crystals, the number of atoms would be much larger.) After one half-life, 8 isotopes have decayed, so the crystal now contains 8 parent and 8 daughter isotopes. After a second half-life, 4 of the remaining parent isotopes have decayed, so the crystal contains 4 parent and 12 daughter isotopes. And after a third half-life, 2 more parent isotopes have  decayed, so the crystal contains 2 parent and 14 daughter isotopes. For a given decay reaction, the half-life is a constant.

Isotopic Dating 

Techniques Since radioactive decay proceeds at a known rate, like the tick-tock of a clock, it provides a basis for telling time. In other words, because an element’s half-life is a constant, we can calculate the age of a mineral by measuring the ratio of parent to daughter isotopes in the mineral. In practice, how can we obtain an isotopic date? First, we must find the right kind of elements to work with. Although there are many different pairs of parent and daughter isotopes among the known radioactive elements, only a few have long enough half-lives, and occur in sufficient abundance in minerals, to be useful for isotopic dating. 

Table above lists particularly useful elements. Each radioactive element has its own half-life. (Note that carbon dating is not used for dating rocks because appropriate carbon isotopes occur only in organisms and radioactive carbon has a very short half-life). Second, we must identify the right kind of minerals to work with. Not all minerals contain radioactive elements, but fortunately some fairly common minerals do. Once we have found the right kind of minerals, we can set to work using the following steps. 

• Collecting the rocks: We need to find un-weathered rocks for dating, for the chemical reactions that happen during weathering may lead to the loss of some isotopes. 
• Separating the minerals: The rocks are crushed, and the appropriate minerals are separated from the debris. 
• Extracting parent and daughter isotopes: To separate out the parent and daughter isotopes from minerals, we can use several techniques, including dissolving the minerals in acid or evaporating portions of them with a laser. 
• Analyzing the parent-daughter ratio: Once we have a sample of appropriate atoms, we pass them through a mass spectrometer, an instrument that uses a strong magnet to separate isotopes from one another according to their respective weights (figure below). The instrument can count the number of atoms of specific isotopes separately. 

At the end of the laboratory process, we can define the ratio of parent to daughter isotopes in a mineral, and from this ratio calculate the age of the mineral. Needless to say, the description of the procedure here has been simplified in reality, obtaining an isotopic date is time-consuming and expensive and requires complex calculations.

What Does an Isotopic Date Mean? 

At high temperatures, atoms in a crystal lattice vibrate so rapidly that chemical bonds can break and reattach relatively easily. As a consequence, isotopes can escape from or move into crystals, so parent-daughter ratios are meaningless. Because isotopic dating is based on the parent-daughter ratio, the “isotopic clock” starts only when crystals become cool enough for isotopes to be locked into the lattice. The temperature below which isotopes are no longer free to move is called the closure temperature of a mineral. When we specify an isotopic date for a mineral, we are defining the time at which the mineral cooled below its closure temperature. With the concept of closure temperature in mind, we can interpret the meaning of isotopic dates. In the case of igneous rocks, isotopic dating tells you when a magma or lava cooled to form a solid, cool igneous rock. In the case of metamorphic rocks, an isotopic date tells you when a rock cooled from a metamorphic temperature above the closure temperature to a temperature below. Not all minerals have the same closure temperature, so different minerals in a rock that cools very slowly will yield different dates. Can we isotopically date a clastic sedimentary rock directly? No. If we date minerals in a sedimentary rock, we determine only when these minerals first crystallized as part of an igneous or metamorphic rock, not the time when the minerals were deposited as sediment nor the time when the sediment lithified to form a sedimentary rock. For example, if we date the feldspar grains contained within a granite pebble in a conglomerate, we’re dating the time the granite cooled below feldspar’s closure temperature, not the time the pebble was deposited by a stream.

Other Methods of Determining Numerical Age

The rate of tree growth depends on the season. During the spring, trees grow rapidly and produce lighter, less-dense wood, but during the winter trees grow slowly or not at all, and produce darker, denser wood. Thus, wood contains recognizable annual growth rings. Such tree rings provide a basis for determining age. If you've ever wondered how old a tree that’s just been cut down might be, just look at the stump and count the rings. Notably, by correlating clusters of distinctive rings in the older parts of living trees with comparable clusters of rings in dead logs, scientists can extend the tree-ring record back for many thousands of years, allowing geologists to track climate changes back into prehistory. Seasonal changes also affect rates of such phenomena as shell growth, snow accumulation, clastic sediment deposition, chemical sediment precipitation, and production of organic material. Geologists have learned to use growth rings in shells, as well as rhythmic layering in sediments and in glacial ice (figure above a–c), to date events numerically back through recent Earth history.

How do Earthquakes causes damage?

Damages from Earthquakes

An area ravaged by a major earthquake is a heartbreaking sight. The terror and sorrow etched on the faces of survivors mirror the inconceivable destruction. This destruction comes as a result of many processes.

Ground Shaking and Displacement 

An earthquake starts suddenly and may last from a few seconds to a few minutes. Different kinds of earthquake waves cause different kinds of ground motion (figure above). The nature and severity of the shaking at a given location depend on four factors: 

1. the magnitude of the earthquake, because larger magnitude events release more energy; 
2. the distance from the focus, because earthquake energy decreases as waves pass through the Earth; 
3. the nature of the substrate at the location (that is, the character and thickness of different materials beneath the ground surface) because earthquake waves tend to be amplified in weaker substrate; and 
4. the “frequency” of the earthquake waves (where frequency equals the number of waves that pass a point in a specified interval of time). 

If you’re out in an open field during an earthquake, ground motion alone won’t kill you, you may be knocked off your feet and bounced around a bit, but your body is too flexible to break. Buildings and bridges aren't so lucky (figure above a-d). When earthquake waves pass, they sway, twist back and forth, or lurch up and down, depending on the type of wave motion. As a result, connectors between the frame and facade of a building may separate, so the facade crashes to the ground. The flexing of walls shatters windows and makes roofs collapse. Floors or bridge decks may rise up and slam down on the columns that support them, thereby crushing the columns. Some buildings collapse with their floors piling on top of one another like pancakes in a stack, some crumble into fragments, and some simply tip over. The majority of  earthquake-related deaths and injuries happen when people are hit by debris or are crushed beneath falling walls or roofs. Aftershocks worsen the problem, because they may topple already weakened buildings, trapping rescuers. During earthquakes, roads, rail lines, and pipelines may also buckle and rupture. If a building, fence, road, pipeline, or rail line straddles a fault, slip on the fault can crack the structure and separate it into two pieces.

Landslides 

The shaking of an earthquake can cause ground on steep slopes or ground underlain by weak sediment to give way. This movement results in a landslide, the tumbling and flow of soil and rock down-slope. Earthquake triggered landslides occur commonly along the coast of  California where expensive homes perch on steep cliffs looking out over the Pacific. When the cliffs collapse, the homes may tumble to the beach below (figure above a and b). Such events lead to the misperception that “California will someday fall into the sea.” Although small portions of the coastline do collapse, the state as a whole remains firmly attached to the continent, despite what  Hollywood  scriptwriters say.

Sediment Liquefaction 

In 1964, an MW 7.5 earthquake struck Niigata, Japan. A portion of the city had been built on land underlain by wet sand. During the ground shaking, foundations of over 15,000 buildings sank into their substrate, causing walls and roofs to crack. Several four-story-high buildings in a newly built apartment complex tipped over (figure above c). The same year, on Good Friday, an MW  9.2 earthquake devastated southern Alaska. In the Turnagain Heights neighbourhood of Anchorage, the event led to catastrophe. The neighbourhood was built on a small terrace of uplifted sediment. The edge of the terrace was a 20-m high escarpment that dropped down to Cook Inlet, a bay of the Pacific Ocean. As the ground shaking began, a layer of wet clay beneath the development turned into mud, and when this happened, the overlying layers of sediment, along with the houses built on top of them, slid seaward. In the process, the layers broke into separate blocks that tilted, turning the landscape into a chaotic jumble, and resulting in the destruction of the neighbourhood (figure above d). 

In 2011, an earthquake in Christchurch, New Zealand, caused sand to erupt and produce small, cone-shaped mounds on the ground surface  (figure above a). The transfer of sand from underground onto the surface led to formation of depressions large enough to  swallow cars (figure above b). All of these examples are manifestations of a phenomenon called sediment liquefaction. During liquefaction, pressure in the water filling the pores between grains in wet sand push the grains apart so that they become surrounded by water and no longer rest against each other. In wet clay, shaking breaks the weak electrostatic charges that hold clay flakes together, so what had been a gel-like, stable mass becomes slippery mud. As the material above the liquefied sediment settles downward, pressure can squeeze the sand upward and out onto the ground surface. The resulting cone-shaped mounds are variously known as sand volcanoes, sand boils, or sand blows. The settling of sedimentary layers down into a liquefied layer can also disrupt bedding and can lead to formation of open fissures of the land surface (figure above c).

Fire 

The shaking during an earthquake can make lamps, stoves, or candles with open flames tip over, and it may break wires or topple power lines, generating sparks. As a consequence, areas already turned to rubble, and even areas not so badly damaged may be consumed by fire. Ruptured gas pipelines and oil tanks feed the flames, sending columns of fire erupting skyward (figure above). Fire fighters might not even be able to reach the fires, because the doors to the fire house won’t open or rubble blocks the streets. Moreover, fire fighters may find themselves without water, for ground shaking and landslides damage water lines. Once a fire starts to spread, it can become an unstoppable inferno. Most of the destruction of the 1906 San Francisco earthquake, in fact, resulted from fire. For three days, the blaze spread through the city until fire fighters contained it by blasting a fire break. By then, 500 blocks of structures had turned to ash, causing 20 times as much financial loss as the shaking itself. When a large earthquake hit Tokyo in 1923, fires set by cooking stoves spread quickly through the wood-and-paper buildings, creating an inferno a “fire storm” that heated the air above the city. As hot air rose, cool air rushed in, creating wind gusts of over 100 mph, which stoked the blaze and incinerated 120,000 people.

Tsunamis 

The azure waters and palm fringed islands of the Indian Ocean’s east coast hide one of the most seismically active plate boundaries on Earth, the Sunda Trench. Along this convergent boundary, the Indian Ocean floor subducts at about 4 cm per year, leading to slip on large thrust faults. Just before 8:00 a.m. on December 26, 2004, the crust above a 1,300-km long by 100-km-wide portion of one of these faults lurched westward by as much as 15 m. The break started at the hypocentre and then propagated north at 2.8 km/s; thus, the rupturing took 9 minutes. This slip triggered a great earthquake (MW 9.3) and pushed the sea floor up by tens of centimetres. The rise of the sea floor, in turn, shoved up the overlying water. Because the area that rose was so broad, the volume of displaced water was immense. As a consequence, a tragedy of an unimaginable extent was about to unfold. Water from above the up-thrust sea floor began moving outward from above the fault zone, a process that generated a series of giant waves travelling at speeds of about 800 km per hour (500  mph) almost the speed of a jet plane (figure above a). Geologists now use the term tsunami for a wave produced by displacement of the sea floor. The displacement can be due to an earthquake, submarine landslide, or volcanic explosion. Tsunami is a Japanese word that translates literally as harbour wave, an apt name because tsunamis can be particularly damaging to harbour towns. In older literature such waves were called “tidal waves,” because when one arrives, water rises as if a tide were coming in, but in fact the waves have nothing to do with daily tidal cycles. Regardless of cause, tsunamis are very different from familiar, wind-driven storm waves. Large wind-driven waves can reach heights of 10 to 30 meters in the open ocean. But even such monsters have wavelengths of only tens of meters, and thus contain a relatively small volume of water. In contrast, although a tsunami in deep water may cause a rise in sea level of at most only a few tens of centimetres a ship crossing one wouldn't even notice tsunamis have wavelengths of tens to hundreds of kilometres and an individual wave can be several kilometres wide, as measured perpendicular to the wave front. Thus, the wave involves a huge volume of water. In simpler terms, we can think of the width of a tsunami, in map view, as being more than 100 times the width of a wind-driven wave. Because of this difference, a storm wave and a tsunami have very different effects when they strike the shore. When a wave approaches the shore, friction between the base of the wave and the sea floor slows the bottom of the wave, so the back of the wave catches up to the front, and the added volume of water builds the wave higher (figure above b). The top of the wave may fall over the front of the wave and cause a breaker. In the case of a wind-driven wave, the breaker may be tall when it washes onto the beach, but because the wave doesn't contain much water, the wave runs out of water and friction slows it to a stop on the beach. Then, gravity causes the water to spill back seaward. In the case of a tsunami, the wave is so wide that, as friction slows the wave, it builds into a “plateau” of water that can be tens of meters high, many kilometres wide, and hundreds of kilometres long. Thus, when a tsunami reaches shore, it contains so much water that it crosses the beach and, if the land is low-lying, just keeps on going, eventually covering a huge area (figure above c). 

Tsunami damage can be catastrophic. The December 2004 waves struck Banda Aceh, a city at the north end of the island of Sumatra, on a beautiful, cloudless day (figure above a). First, the sea receded much farther than anyone had ever seen, exposing large areas of reefs that normally remained submerged even at low tide. People walked out onto the exposed reefs in wonder. But then, with a rumble that grew to a roar, a wall of frothing water began to build in the distance and approach land (figure above b). Puzzled bathers first watched, then ran inland in panic when the threat became clear. As the tsunami approached shore, friction with the sea floor had slowed it to less than 30 km an hour, but it still moved faster than people could run. In places, the wave front reached heights of 15 to 30 m (45 to 100 feet) as it slammed into Banda Aceh (figure above c). The impact of the water ripped boats from their moorings, snapped trees, battered buildings into rubble, and tossed cars and trucks like toys. And the water just kept coming, eventually flooding low-lying land up to 7 km inland (figure above d). It drenched forests and fields with salt water (deadly to plants) and buried fields and streets with up to a meter of sand and mud. When the water level finally returned to normal, a jumble of flotsam, as well as the bodies of unfortunate victims, were dragged out to sea and drifted away. Geologists refer to the tsunami that struck Banda Aceh as a near-field (or local) tsunami, because of its proximity to the earthquake. But the horror of Banda Aceh was merely a preamble to the devastation that would soon visit other stretches of Indian Ocean coast. Far-field (or distant) tsunamis crossed the ocean and struck Sri Lanka 2.5 hours after the earthquake, the coast of India half an hour after that, and the coast of Africa, on the west side of the Indian Ocean, 5.5 hours after the earthquake. In the end, more than 230,000 people died that day. The tsunami that struck Japan soon after the 2011 Tohoku earthquake was vividly captured in high-definition video that was seen throughout the world, generating a new level of international awareness. Though much of the coast was fringed by seawalls, they proved to be a minor impediment to the advance of the wave, which, in places, was 10 to 30 m high when it reached shore. Racing inland the wave erased whole towns, submerging airports and fields. As the wave picked up dirt and debris, it became a viscous slurry, moving with such force that nothing could withstand its impact. 

The devastation of coastal towns was so complete that they looked as though they had been struck by nuclear bombs  (figure above a). But the catastrophe was not over. The wave had also hit a nuclear power plant. Though the plant had withstood ground shaking and had automatically shut down, its radioactive core still needed to be cooled by water in order to remain safe. The tsunami not only destroyed power lines, cutting the plant off from the electrical grid, but it also eliminated backup diesel generators and cut water lines. Thus, cooling pumps stopped functioning. Eventually, water surrounding the heat producing radioactive core of the reactors, as well as the water cooling spent fuel, boiled away. Some of the water separated into hydrogen and oxygen gas, which exploded, and ultimately, the integrity of the nuclear plant was breached so that radioactivity entered the environment (figure above b). Because tsunamis are so dangerous, predicting their arrival can save thousands of lives. A tsunami warning centre in Hawaii keeps track of earthquakes around the Pacific and uses data relayed from tide gauges and sea-floor pressure gauges to determine whether a particular earthquake has generated a tsunami. If observers detect a tsunami, they flash warnings to authorities around the Pacific. 

Disease 

Once the ground shaking and fires have stopped, disease may still threaten lives in an earthquake damaged region. Earthquakes cut water and sewer lines, destroying clean-water supplies and exposing the public to bacteria, and they cut transportation lines, preventing food and medicine from reaching the area. The severity of such problems depends on the ability of emergency services to cope. The lack of sufficient clean water after the 2010 Haiti earthquake led to a cholera epidemic later that year.

Can we predict Earthquakes?

Can seismologists predict earthquakes? 

The answer depends on the time frame of the prediction. With our present understanding of the distribution of seismic zones and the frequency at which earthquakes occur, we can make long-term predictions (on the time scale of decades to centuries). For example, with some certainty, we can say that a major earthquake will rattle Istanbul during the next 100 years, and that a major earthquake probably won’t strike central Canada during the next 10 years. But despite extensive research, seismologists cannot make accurate short-term predictions (on the time scale of hours to weeks or even years). Thus we cannot say, for example, that an earthquake will happen in Montreal at 2:43 P.M. on January 17. In this section, we look at the scientific basis of both long- and short-term predictions and consider the consequences of a prediction. Seismologists refer to studies leading to predictions as seismic-risk, or  seismic-hazard assessment. 

Long-Term Predictions 

A long-term prediction estimates the probability, or likelihood that an earthquake will happen during a specified time range. For example, a seismologist may say, “The probability of a major earthquake occurring in the next 20 years in this state is 20%.” This sentence implies that there’s a 1-in-5 chance that the earthquake will happen before 20 years have passed. Urban planners and civil engineers can use long-term predictions to help create building codes for a region codes requiring stronger, more expensive buildings make sense for regions with greater seismic risk. They may also use predictions to determine whether it is reasonably safe to build vulnerable structures such as nuclear power plants, hospitals, or dams in a given region. Seismologists base long-term earthquake predictions on two pieces of information: the identification of seismic zones and the recurrence interval (the average time between successive events). 

To identify a seismic zone, seismologists produce a map showing the epicentres of earthquakes that have happened  during a set period of time (say, 30 years). Clusters or belts of epicentres define the seismic zone. The basic premise of long-term earthquake prediction can be stated as follows: a region in which there have been many earthquakes in the past will be more likely to experience earthquakes in the future. Seismic zones, therefore, are regions of greater seismic risk. This doesn't mean that a disastrous earthquake can’t happen far from a seismic zone they can and do but the probability that an event will happen in a given time window is less. To determine the recurrence interval for large earthquakes within a given seismic zone, seismologists must determine when large earthquakes happened there in the past. Since the historical record does not provide information far enough back in time, they study geologic evidence for great earthquakes. For example, recognition of a fresh, unweathered fault scarp or trace may indicate that faulting affected an area relatively recently. A trench cut into sedimentary strata near a fault may reveal layers of sand volcanoes and disrupted bedding in the stratigraphic record. Each layer, whose age can be determined by using radiocarbon dating of plant fragments, records the time of an earthquake (figure above). By calculating the number of years between successive events and taking the average, seismologists obtain the recurrence interval. Note that a recurrence interval does not specify the exact number of years between events, only the average number. Since stress builds up over time on a fault, the probability that an earthquake will happen in any given year probably increases as time passes. 

Information on a recurrence interval allows seismologists to refine regional maps illustrating seismic risk (figure above a and b).

Short-Term Predictions 

Short-term predictions, specifying that an earthquake will happen on a given date or within a time window of days to years, are not and may never be reliable. Seismologists have considered, and discounted as unreliable, many supposed bases for short-term prediction. For example, a swarm of fore-shocks may indicate that rock is beginning to crack in advance of a main-shock, but such swarms can be identified only in hindsight. Precise surveys show that the surface of the ground may warp slightly prior to an earthquake, but no one can determine how much warping will take place before an earthquake will happen. Prediction studies focused on measuring changes in water levels in wells, radon gas in spring water, electrical signals emitted by minerals, or agitation of animals have met with similar skepticism. The concept of a short-term prediction should not be confused with the concept of an earthquake early warning system. An early warning system works as follows. When an earthquake happens, the seismic waves it produces start travelling through the Earth. Seismic stations closer to the epicentre may detect an earthquake before the seismic waves have had time to reach populated areas farther from the epicentre. The instant that seismic stations detect the earthquake, a computer approximates the epicentre location, then sends a signal to a control centre, which automatically sends out emergency signals to areas that might be affected. The signals shut down gas pipelines, trains, nuclear reactors, power lines, and other vulnerable infrastructure. The signal also sets off sirens and alerts broadcasters to send out warnings on radio, TV, and cell-phone networks to warn people that an earthquake is about to begin. Unless the focus is directly under the city, the warning may precede the arrival of the first earthquake waves by several seconds, not a lot of time, but hopefully enough to prevent some infrastructure damage and perhaps enough for people to seek a safer location.

Consequences and Causes of Metamorphism

What Is a Metamorphic Rock? 

If someone were to put a rock on a table in front of you, how would you know that it is metamorphic? First, metamorphic rocks can possess metamorphic minerals, new minerals that grow in place within the solid rock only under metamorphic temperatures and pressures. In fact, metamorphism can produce a group of minerals that together make up what geologists call a “metamorphic mineral assemblage.” And second, metamorphic rocks can have metamorphic texture defined by distinctive arrangements of mineral grains not found in other rock types. Commonly, the texture results in metamorphic foliation, due to the parallel alignment of platy minerals (such as mica) and/ or the presence of alternating light-coloured and dark-coloured layers. When metamorphic minerals and/or textures develop, a metamorphic rock becomes as different from its protolith as a butterfly is from a caterpillar. For example, metamorphism of red shale can yield a metamorphic rock consisting of aligned mica flakes and brilliant garnet crystals (a in figure above), and metamorphism of a limestone composed of cemented-together fossil fragments can yield a metamorphic rock consisting of large interlocking crystals of calcite (b in figure above). The process of forming metamorphic minerals and textures takes place very slowly it may take thousands to millions of years and it involves several processes, which sometimes occur alone and sometimes together. The most common processes are: 

• Recrystallization, which changes the shape and size of grains without changing the identity of the mineral making up the grains (a in figure above). 
• Phase change, which transforms one mineral into another mineral with the same composition but a different crystal structure. On an atomic scale, phase change involves the rearrangement of atoms. 
• Metamorphic reaction, or neocrystallization (from the Greek neos, for new), which results in growth of new mineral crystals that differ from those of the protolith (b in figure above). During neocrystallization, chemical reactions digest minerals of the protolith to produce new minerals of the metamorphic rock. 
• Pressure solution, which happens when a wet rock is squeezed more strongly in one direction than in others. Mineral grains dissolve where their surfaces are pressed against other grains, producing ions that migrate through the water to precipitate elsewhere (c in figure above). 
• Plastic deformation, which happens when a rock is squeezed or sheared at elevated temperatures and pressures. Under these conditions, grains behave like soft plastic and change shape without breaking (d in figure above). 

Caterpillars undergo metamorphosis because of hormonal changes in their bodies. Rocks undergo metamorphism when they are subjected to heat, pressure, compression and shear, and/or very hot water. Let’s now consider the details of how these agents of metamorphism operate.

Metamorphism Due to Heating 

When you heat cake batter, the batter transforms into a new material cake. Similarly, when you heat a rock, its ingredients transform into a new material metamorphic rock. Why? Think about what happens to atoms in a mineral grain as the grain warms. Heat causes the atoms to vibrate rapidly, stretching and bending chemical bonds that lock atoms to their neighbours. If bonds stretch too far and break, atoms detach from their original neighbours, move slightly, and form new bonds with other atoms. Repetition of this process leads to rearrangement of atoms within grains, or to migration of atoms into and out of grains, a process called solid-state diffusion. As a consequence, recrystallization and/or neo-crystallization take place, enabling a metamorphic mineral  assemblage to grow in solid rock. Metamorphism takes place at temperatures between those at which diagenesis occurs and those that cause melting. Roughly speaking, this means that most metamorphic rocks you find in outcrops on continents formed at temperatures of between 250C and 850C.

Metamorphism Due to Pressure 

As you swim underwater in a swimming pool, water squeezes against you equally from all sides in other words, your body feels pressure. Pressure can cause a material to collapse inward. For example, if you pull an air-filled balloon down to a depth of 10 m in a lake, the balloon becomes significantly smaller. Pressure can have the same effect on minerals. Near the Earth’s surface, minerals with relatively open crystal structures can be stable. However, if you subject these minerals to extreme pressure, the atoms pack more closely together and denser minerals tend to form. Such transformations involve phase changes and/or neo-crystallization.

Changing Both Pressure and Temperature 

So far, we've considered changes in pressure and temperature as separate phenomena. But in the Earth, pressure and temperature change together with increasing depth. For example, at a depth of 8 km, temperature in the crust reaches about 200C and pressure reaches about 2.3 kbar. If a rock slowly becomes buried to a depth of 20 km, as can happen during mountain building, temperature in the rock increases to more than 500C, and pressure to 5.5 kbar. Experiments and calculations show that the “stability” of certain minerals (the ability of a mineral to form and survive) depends on both pressure and temperature. When pressure and temperature increase, the original mineral assemblage in a rock becomes unstable, and a new assemblage forms out of minerals that are stable. Thus, a metamorphic rock formed at 8 km does not contain the same minerals as one formed at 20 km.

Compression, Shear, and Development  of Preferred Orientation 

Imagine that you have just built a house of cards and, being in a destructive mood, you step on it. The structure collapses because the downward push you apply with your foot exceeds the push provided by air in other directions. We can say that we have subjected the cards to compression (a in figure above). Compression flattens a material (b in figure above). Shear, in contrast, moves one part of a material sideways, relative to another. If, for example, you place a deck of cards on a table, then set your hand on top of the deck and move your hand parallel to the table, you shear the deck (c in figure above). When rocks are subjected to compression and shear at elevated temperatures and pressures, they can change shape without breaking. As it changes shape, the internal texture of a rock also changes. For example, platy (pancake-shaped) grains become parallel to one another, and elongate (cigar shaped) grains align in the same direction. Both platy and elongate grains are inequant grains, meaning that the dimension of a grain is not the same in all directions; in contrast, equant grains have roughly the same dimensions in all directions (d in figure above). The alignment of inequant minerals in a rock results in a preferred orientation (e in figure above).

The Role of Hydrothermal Fluids 

Metamorphic reactions commonly take place in the presence of hydrothermal fluids (very hot-water solutions). Where does the water in hydrothermal fluids come from? Some of it was originally bonded to minerals in the protolith, for metamorphic reactions can release such water into its surroundings. Some of it may seep up into the protolith from a nearby igneous intrusion, or down from overlying groundwater reservoirs. Notably, under extremely high pressures and temperatures, the water of hydrothermal fluids is in neither gas nor liquid state, but rather is in a “supercritical” state, meaning that it has characteristics of both gas and liquid. Such hydrothermal fluids chemically react with rock; they accelerate metamorphic reactions, because atoms involved in the reactions can migrate faster through a fluid than they can through a solid, and hydrothermal fluids provide water that can be absorbed by minerals during metamorphic reactions. Finally, fluids passing through a rock may pick up some dissolved ions and drop off others, as a bus picks up and drops off passengers, and thus can change the overall chemical composition of a rock during metamorphism. The process of changing a rock’s chemical composition by reactions with hydrothermal fluids is called metasomatism.

Recognizing Depositional Environments

How Do We Recognize Depositional Environments? 

Geologists refer to the conditions in which sediment was deposited as the depositional environment. Examples include beach, glacial, and river environments. To identify depositional environments, geologists, like crime scene investigators, look for clues. Detectives may seek fingerprints and bloodstains to identify a culprit. Geologists examine grain size, composition, sorting, bed-surface marks, cross bedding, and fossils to identify a depositional environment. Geological clues can tell us if the sediment was deposited by ice, strong currents, waves, or quiet water, and in some cases can provide insight into the climate at the time of deposition. With experience, geologists can examine a succession of beds and determine if it accumulated on a river floodplain, along a beach, in shallow water just offshore, or on the deep ocean floor.

Let’s now explore some examples of different depositional environments and the sediments deposited in them, by imagining that we are taking a journey from the mountains to the sea, examining sediments as we go. We will see that geologists distinguish among three basic categories of depositional environments: terrestrial, coastal, and marine.

Terrestrial (Nonmarine) Sedimentary Environments 

We begin our exploration with terrestrial depositional environments, those that develop inland, far enough away from the shoreline that they are not affected by ocean tides and waves. The sediments settle on dry land, or under and adjacent to freshwater.  

In some settings, oxygen in surface water or groundwater reacts with iron to produce rust-like iron-oxide minerals in terrestrial sediments, which give the sediment an overall reddish hue. Strata with this hue are informally called redbeds.

Glacial environments 

High in the mountains, where it’s so cold that more snow collects in the winter than melts away,  glaciers rivers or sheets of ice develop and slowly flow. Because ice is a solid, it can move sediment of any size. So as a glacier moves down a valley in the mountains, it carries along all the sediment that falls on its surface from adjacent cliffs or gets plucked from the ground at its base or sides. At the end of the glacier, where the ice finally melts away, the sediment that had been in or on the ice accumulates as “glacial till” (a in figure above). Till is unsorted and unstratified it contains clasts ranging from clay size to boulder size all mixed together.

Mountain stream environments

As we walk down beyond the end of the glacier, we enter a realm where turbulent streams rush downslope in steep-sided valleys. This fast-moving water has the power to carry large clasts; in fact, during floods, boulders and cobbles can tumble down the stream bed. Between floods, when water flow slows, the largest clasts settle out to form gravel and boulder beds, while the stream carries finer sediments like sand and mud farther downstream (b in figure above). Sedimentary deposits of a mountain stream would, therefore, include breccia and  conglomerate.

Alluvial-fan environments

Our journey now takes us to the mountain front, where the fast-moving stream empties onto a plain. In arid regions, where there is not enough water for the stream to flow continuously, the stream deposits its load of sediment near the mountain front, producing a wedge-shaped apron of gravel and sand called an alluvial fan  (c in figure above). Deposition takes place here because when the stream pours from a canyon mouth and spreads out over a broader region, friction with the ground causes the water to slow down, and slow-moving water does not have the power to move coarse sediment. The sand here still contains feldspar grains, for these have not yet weathered into clay. Alluvial-fan sediments become arkose and conglomerate.

Sand-dune environments

If the climate is very dry, few plants can grow and the ground surface lies exposed. Strong winds can move dust and sand. The dust gets carried away, and the resulting well-sorted sand can accumulate in dunes. Thus, thick layers of well-sorted sandstone, in which we can find large cross beds, are relicts of desert sand-dune environments (d in figure above).

River (fluvial) environments

In climates where streams flow, we find several distinctive depositional environments. Rivers transport gravel, sand, silt, and mud. The coarser sediments tumble along the bed in the river’s channel and collect in cross-bedded, rippled layers while the finer sediments drift along, suspended in the water. This fine sediment settles out along the banks of the river, or on the floodplain, the flat land on either side of the river that is covered with water only during floods. On the floodplain, mud layers dry out between floods, leading to the formation of mud cracks. River sediments lithify to form sandstone, siltstone, and shale. Typically, the coarser sediments of channels are surrounded by layers of fine-grained floodplain deposits, so in cross section, the channel has a lens-like shape (e in figure above). Geologists commonly refer to river deposits as fluvial sediments, from the Latin word fluvius, for river.

Lake environments

In temperate climates, where water remains at the surface throughout the year, lakes form. In lakes, the relatively quiet water can’t move coarse sediment; any coarse sediment brought into the lake by a stream settles out at the stream’s outlet. Only fine clay makes it out into the centre of the lake, where it settles to form mud on the lake bed. Thus, lake sediments typically consist of finely  laminated shale (f in figure above). 

At the mouths of streams that empty into lakes, small deltas may form. A delta is a wedge of sediment that accumulates where moving water enters standing water. Deltas were so named because the map shape of some deltas resembles the Greek letter delta ($), as we discuss further in Chapter 14. In 1885, an American geologist named G. K. Gilbert showed that such deltas contain three components (figure above): topset beds composed of gravel, foreset beds of gravel and sand, and silty bottomset beds.

Coastal and Marine Environments 

Along the seashore, a variety of distinct coastal environments occur; the character of each reflects the nature of the sediment supply and the climate. Marine environments start at the high-tide line and extend offshore, to include the deep ocean floor. The type of sediment deposited at a location depends on the climate, water depth, and whether or not clastic grains are available.

Marine delta deposits

After following the river downstream for a long distance, we reach its mouth, where it empties into the sea. Here, the river builds a delta of sediment out into the sea. River water stops flowing when it enters the sea, so sediment settles out. Large deltas are much more complex than the lake examples that Gilbert studied, for they include many different sedimentary environments including swamps, channels, floodplains, and submarine slopes. Sea-level changes may cause the positions of the different environments to move with time. Thus, deposits of an ocean-margin delta produce a great variety of sedimentary rock types (a in figure above).

Coastal beach sands

Now we leave the delta and wander along the coast. Oceanic currents transport sand along the coastline. The sand washes back and forth in the surf, so it becomes well sorted (waves winnow out silt and clay) and well rounded, and because of the back-and-forth movement of ocean water over the sand, the sand surface may become rippled (b in figure above). Thus, if you find well-sorted, medium grained sandstone, perhaps with ripple marks, you may be looking at the remnants of a beach environment.

Shallow-marine clastic deposits

From the beach, we proceed offshore. In deeper water, where wave energy does not stir the sea floor, finer sediment can accumulate. Because the water here may be only meters to a few tens of meters deep, geologists refer to this depositional setting as a shallow-marine environment. Clastic sedimentary layers that accumulate in this environment tend to be fine-grained, well-sorted, well rounded silt, and they are inhabited by a great variety of organisms such as mollusks and worms. Thus, if you see beds of siltstone and mudstone containing marine fossils, you may be looking at shallow-marine clastic deposits.

Shallow-water carbonate environments

In shallow marine settings relatively free of clastic sediment, warm, clear, nutrient-rich water hosts an abundance of organisms. Their shells, which consist of carbonate minerals, make up most of the sediment that accumulates (a and b in figure above). The nature of carbonate sediment depends on the water depth. Beaches collect sand composed of shell fragments; lagoons (protected bodies of quiet water) are sites where carbonate mud accumulates; and reefs consist of coral and coral debris. Farther offshore of a reef, we can find a sloping apron of reef fragments. Shallow-water carbonate environments transform into various kinds of limestone.

Deep-marine deposits

We conclude our journey by sailing far offshore. Along the transition between coastal regions and the deep ocean, turbidity currents deposit graded beds. In the deep-ocean realm, only fine clay and plankton provide a source for sediment. The clay eventually settles out onto the deep-sea floor, forming deposits of finely laminated mudstones, and plankton shells settle to form chalk (from calcite shells; a and b in figure above) or chert (from siliceous shells). Thus, deposits of mudstone, chalk, or bedded chert indicate a deepmarine origin.

Sedimentary structures

Sedimentary structures

Geologists use the term sedimentary structure for the layering of sedimentary rocks, for surface features on layers formed during deposition, and for the arrangement of grains within layers. Here, we examine some of the more important types.

Bedding and Stratification 

Let’s start by introducing the jargon for discussing sedimentary layers. A single layer of sediment or sedimentary rock with a recognizable top and bottom is called a bed; the boundary between two beds is a bedding plane; several beds together constitute strata (singular stratum, from the Latin stratum, meaning pavement); and the overall arrangement of sediment into a sequence of beds is bedding, or stratification. From the word strata, we derive other words, such as stratigrapher (a geologist who specializes in studying strata) and stratigraphy (the study of the record of Earth history preserved in strata). In some outcrops, stratification can be quite subtle. But commonly, successive beds have different colours, textures, and resistance to erosion, so bedding gives outcrops a striped appearance (a in figure above).

Why does bedding form? To find the answer, we need to think about how sediment accumulates. Changes in the climate, water depth, current velocity, or the sediment source control the type of sediment deposited at a location at a given time. For example, on a normal day a slow- moving river may carry only silt, which collects on the riverbed (b in figure above). During a flood, the river flows faster and carries sand and pebbles, so a layer of sandy gravel forms over the silt layer. Then, when the flooding stops, more silt buries the gravel. If this  succession of sediments become lithified and exposed for you to see, they appear as alternating beds of siltstone and sandy conglomerate. During geologic time, long-term changes in a depositional environment can take place. Thus, a given sequence of strata may differ markedly from sequences of strata above or below. 

A sequence of strata that is distinctive enough to be traced as a unit across a fairly large region is called a stratigraphic formation, or simply a formation (a in figure above). For example, a region may contain a succession of alternating sandstone and shale beds deposited by rivers, overlain by beds of marine limestone deposited later when the region was submerged by the sea. A stratigrapher might identify the sequence of sandstone and shale beds as one formation and the sequence of limestone beds as another. Formations are often named after the locality where they were first found and studied. A map that portrays the distribution of stratigraphic formations is called a geologic map (b in figure above).

Ripple Marks, Dunes, and Cross Bedding: Consequences of Deposition in a Current 

Many clastic sediments accumulate in moving fluids (wind, rivers, or waves). Fascinating sedimentary structures develop at the interface between the sediment and the fluid. These structures are called bedforms. Bedforms that develop at a given location reflect such factors as the velocity of the flow and the size of the clasts. Though there are many types of bedforms, we’ll focus on only two ripple marks and dunes. The growth of both produces cross bedding, a special type of lamination within beds. Ripple marks are relatively small (generally no more than a few centimetres high), elongated ridges that form on a bed surface at right angles to the direction of current flow. You can find ripples on modern beaches and preserved on bedding planes of ancient rocks (figure above). Dunes are relatively large, elongate ridges built of sediment transported by a current. In effect, dunes are “mega-ripples.” Dunes on the bed of a stream may be tens of centimeters high, and wind-formed dunes of deserts may be tens to over 100 meters high. 

If you examine a vertical slice cut into a ripple or dune, you will find distinct internal laminations that are inclined at an angle. Such laminations are called cross beds. To see how cross beds develop, imagine a current of air or water moving uniformly in one direction (a in figure above). The current erodes and picks up clasts from the upstream part of the bedform and deposits them on the downstream or leeward face of the crest. Sediment builds up until gravity causes it to slip down the leeward face. With time, the dune or ripple builds in the downstream direction. The surface of the slip face establishes the shape of the cross beds. Eventually, a new cross-bedded layer builds out over a pre-existing one. The boundary between two successive layers is called the “main bedding,” and the internal curving surfaces within the layer constitute the cross bedding (b and c in figure above).

Turbidity Currents and Graded Beds 

Sediment deposited on a submarine slope tends to be unstable. For example, an earthquake or storm might disturb this sediment and cause it to slip downslope and mix with water to create a murky, turbulent cloud. This cloud is denser than clear water and thus flows downslope like an underwater avalanche (a to c in figure above). We call this moving submarine suspension of sediment a turbidity current. Downslope, the turbidity current slows, and the sediment that it has carried starts to settle out. Larger grains sink faster through a fluid than do finer grains, so the coarsest sediment settles out first. Progressively finer grains accumulate on top, with the finest sediment (clay) settling out last. This process forms a graded bed that is, a layer of sediment in which grain size varies from coarse at the bottom to fine at the top. Geologists refer to a deposit from a turbidity current as a turbidite.

Bed-Surface Markings 

A number of features develop on the surface of a bed as a consequence of events that happen during deposition or soon after, while the sediment layer remains soft. Such bed-surface markings include the following: 

• Mud cracks: If a mud layer dries up after deposition, it cracks into roughly hexagonal plates that typically curl up at their edges. We refer to the openings between the plates as mud cracks (a and b figure above). 
• Scour marks: As currents flow over a sediment surface, they may erode small troughs, called scour marks, parallel to the current flow. 
• Fossils: Fossils are relicts of past life. Some fossils are shell imprints or footprints on a bedding surface. 

Burial and lithification of bed-surface markings can preserve them in the stratigraphic record.

Why Study Sedimentary Structures? 

Sedimentary structures are not just a curiosity, but are important clues that help geologists understand the environment in which sedimentary beds were deposited. For example, the presence of ripple marks and cross bedding indicates that layers were deposited in a current, the presence of mud cracks indicates that the sediment layer was exposed to the air and dried out, and graded beds indicate deposition by turbidity currents. Also, fossil types can tell us whether sediment was deposited along a river or in the deep sea, for different species of organisms live in different environments. In the next section of this chapter, we examine these environments in greater detail.

Classes of sedimentary rocks

Classes of sedimentary rocks

Geologists divide sedimentary rocks into four major classes, based on their mode of origin. 

(1) Clastic sedimentary rock consists of cemented-together clasts, solid fragments and grains broken off of preexisting rocks (the word comes from the Greek klastos, meaning broken); (2) biochemical sedimentary rock consists of shells; (3) organic sedimentary rock consists of carbon-rich relicts of plants or other organisms; and (4) chemical sedimentary rock is made up of minerals that precipitated directly from water solutions. Let’s now look at these major classes in more detail.

Clastic Sedimentary Rocks Formation

Nine hundred years ago, a thriving community of Native Americans inhabited the high plateau of Mesa Verde, Colorado. In hollows beneath huge overhanging ledges, they built multistory stone-block buildings that have survived to this day. Clearly, the blocks are solid and durable they are, after all, rock. But if you were to rub your thumb along one, it would feel gritty, and small grains of quartz would break free and roll under your thumb, for the block consists of quartz sand grains cemented together. Geologists call such rock a sandstone. Sandstone is an example of clastic sedimentary rock. It consists of loose clasts, known as detritus, that have been stuck together to form a solid mass. The clasts can consist of individual minerals (such as grains of quartz or flakes of clay) or of fragments of rock (such as pebbles of granite). Formation of sediment and its transformation into clastic sedimentary rock takes place via the following five steps.

• Weathering: Detritus forms by disintegration of bedrock into separate grains due to physical and chemical weathering. 
• Erosion: Erosion refers to the combination of processes that separate rock or regolith (surface debris) from its substrate. Erosion involves abrasion, falling, plucking, scouring, and dissolution, and is caused by moving air, water, or ice. 
• Transportation: Gravity, wind, water, or ice carry sediment. The ability of a medium to carry sediment depends on its viscosity and velocity. Solid ice can transport sediment of any size, regardless of how slowly the ice moves. Very fast-moving, turbulent water can transport coarse fragments (cobbles and boulders), moderately fast-moving water can carry only sand and gravel, and slow-moving water carries only silt and clay. Strong winds can move sand and dust, but gentle breezes carry only dust. 
• Deposition: Deposition is the process by which sediment settles out of the transporting medium. Sediment settles out of wind or moving water when these fluids slow, because as the velocity decreases, the fluid no longer has the ability to carry sediment. Sediment is deposited by ice when the ice melts. 
• Lithification: Geologists refer to the transformation of loose sediment into solid rock as lithification. The lithification of clastic sediment involves two steps. First, once the sediment has been buried, pressure caused by the weight of overlying material squeezes out water and air that had been trapped between clasts, and clasts press together tightly, a process called compaction. Compacted sediment may then be stuck together to make coherent sedimentary rock by the process of cementation. Cement consists of minerals (commonly quartz or calcite) that precipitate from groundwater and fill the spaces between clasts. 

Classifying clastic sedimentary rocks

Say that you pick up a clastic sedimentary rock and want to describe it sufficiently so that, from your words alone, another person can picture the rock. What characteristics should you mention? Geologists find the following characteristics most useful. 

• Clast size. Size refers to the diameter of fragments or grains making up a rock. Names used for clast size, listed in order from coarsest to finest, are: boulder, cobble, pebble, sand, silt, and clay. 
• Clast composition. Composition refers to the makeup of clasts in sedimentary rock. Clasts may be composed of rock fragments or individual mineral grains. 
• Angularity and sphericity. Angularity indicates the degree to which clasts have smooth or angular corners and edges. Sphericity, in contrast, refers to the degree to which the shape of a clast resembles a sphere. 
• Sorting. Sorting of clasts indicates the degree to which the clasts in a rock are all the same size or include a variety of sizes. Well-sorted sediment consists entirely of sediment of the same size, whereas poorly-sorted sediment contains a mixture of more than one clast size. 
• Character of cement. Not all clastic sedimentary rocks have the same kind of cement. In some, the cement consists  predominantly of quartz, whereas in others, it consists predominantly of calcite. 

With these characteristics in mind, we can distinguish among several common types of clastic sedimentary rocks. This table provides common rock names specialists sometimes use other, more precise names based on more complex classification schemes. The size, angularity, sphericity, and sorting of clasts depends on the medium (water, ice, or wind) that transports the clasts and, in the case of water or wind, on both the velocity of the medium and the distance of transport. The composition of the clasts depends on the composition of rock from which the clasts were derived, and on the degree of chemical weathering that the clasts have undergone. Thus, the type of clasts that accumulate in a sedimentary deposit varies with location. To see how, let’s follow the fate of rock fragments as they gradually move from a cliff face in the mountains via a river to the seashore. Different kinds of sediment develop along the route. Each kind, if buried and lithified, would yield a different type of sedimentary rock.

To start, imagine that some large blocks of granite tumble off a cliff and slam into other blocks already at the bottom. The impact shatters the blocks, producing a pile of angular fragments. If these fragments were to be cemented together before being transported very far, the resulting rock would be breccia (a in above figure). Later, a storm causes the fragments (clasts) to be carried away by a turbulent river. In the water, clasts bang into each other and into the riverbed, a process that shatters them into still smaller pieces and breaks off their sharp edges. As the clasts get carried downstream, they gradually become rounded pebbles and cobbles. When the river water slows, these clasts stop moving and form a mound or bar of gravel. Burial and lithification of these rounded clasts produces conglomerate (b in above figure). If the gravel stays put for a long time, it undergoes chemical weathering. As a consequence, cobbles and pebbles break apart into individual mineral grains, eventually producing a mixture of quartz, feldspar, and clay. Clay is so fine that flowing water easily picks it up and carries it downstream, leaving sand containing a mixture of quartz and some feldspar grains this sediment, if buried and lithified, becomes arkose (c in above figure). Over time, feldspar grains in sand continue to weather into clay so that gradually, during successive events that wash the sediment downstream, the sand loses feldspar and ends up being composed almost entirely of durable quartz grains. Some of the sand may make it to the sea, where waves carry it to beaches, and some may end up in desert dunes. This sediment, when buried and lithified, 

becomes quartz sandstone (d figure below). Meanwhile, silt and clay may accumulate in the flat areas bordering streams, regions called floodplains that become inundated only during floods. And some silt and mud settles in a wedge, called a delta, at the mouth of the river, or in lagoons or mudflats along the shore. The silt, when lithified, becomes siltstone, and the mud, when lithified, becomes shale or mudstone (e figure below).

Biochemical Sedimentary

Rocks The Earth System involves many interactions between living organisms and the physical planet. Numerous organisms have evolved the ability to extract dissolved ions from seawater to make solid shells. When the organisms die, the solid material in their shells survives. This material, when lithified, comprises biochemical sedimentary rock. Geologists recognize several different types of biochemical sedimentary rocks, which we now describe.

Limestone (biochemical)

A snorkeler gliding above a reef sees an incredibly diverse community of coral and algae, around which creatures such as clams, oysters, snails (gastropods), and lampshells (brachiopods) live, and above which plankton float (a figure above). Though they look so different from each other, many of these organisms share an important characteristic: they make solid shells of calcium carbonate (CaCO3). The CaCO3 crystallizes either as calcite or aragonite. (These minerals have the same composition, but different crystal structures.) When the organisms die, the shells remain and may accumulate.  

Rocks formed dominantly from this material are the biochemical version of limestone. Since the principal compound making up limestone is CaCO3, geologists refer to limestone as a type of carbonate rock. Limestone comes in a variety of textures, because the material that forms it accumulates in a variety of ways. For example, limestone can originate from reef builders (such as coral) that grew in place, from shell debris that was broken up and transported, from carbonate mud, or from plankton shells that settled like snow out of water. Because of this variety, geologists distinguish among fossiliferous limestone, consisting of visible fossil shells or shell fragments (b figure above); micrite, consisting of very fine carbonate mud; and chalk, consisting of plankton shells. Experts recognize many other types as well. Typically, limestone is a massive light-gray to darkbluish-gray rock that breaks into chunky blocks it doesn't look much like a pile of shell fragments (c figure above). That’s because several processes change the texture of the rock over time. For example, water passing through the rock not only precipitates cement but also dissolves some carbonate grains and causes new ones to grow.

Chert (biochemical).

If you walk beneath the northern end of the Golden Gate Bridge in California, you will find outcrops of reddish, almost porcelain-like rock occurring in 3- to 15-cm-thick layers (a figure above). Hit it with a hammer, and the rock would crack to form smooth, spoon-shaped (conchoidal) fractures. Geologists call this rock biochemical chert; it’s made from cryptocrystalline quartz (crypto is Greek for hidden), meaning quartz grains that are too small to be seen without the extreme magnification of an electron microscope. The chert beneath the Golden Gate Bridge formed from the shells of silica-secreting plankton that accumulated on the sea floor. Gradually, after burial, the shells dissolved, forming a silica-rich gel. Chert then formed when this gel solidified. 

Organic Sedimentary Rocks 

We've seen how the mineral shells of organisms (CaCO3 or SiO2) can accumulate and lithify to become biochemical sedimentary rocks. What happens to the “guts” of the organisms the cellulose, fat, carbohydrate, protein, and other organic compounds that make up living matter? Commonly, this organic debris gets eaten by other organisms or decays at the Earth’s surface. But in some environments, the organic debris settles along with other sediment and eventually gets buried. When lithified, organic-rich sediment becomes organic sedimentary rock. Since the dawn of the industrial revolution in the early 19th century, coal, one type of organic sedimentary rock, has provided the fuel of modern industry and transportation, for the organic chemicals in the rock yield energy when burned. Coal is a black, combustible rock consisting of over 50 to 90% carbon. The remainder consists of oxygen, nitrogen, hydrogen, sulphur, silica, and minor amounts of other elements. Typically, the carbon in coal occurs in large, complex organic molecules made of many rings note that the carbon does not occur in CaCO3. Coal forms when plant remains have been buried deeply enough and long enough for the material to become compacted and to lose significant amounts of volatiles (hydrogen, water, CO2, and ammonia); as the volatiles seep away, a concentration of carbon remains  (b figure above).

Chemical Sedimentary Rocks 

The colourful terraces, or mounds, that grow around the vents of hot-water springs; the immense layers of salt that underlie the floor of the Mediterranean Sea; the smooth, sharp point of an ancient arrowhead these materials all have something in common. They all consist of rock formed primarily by the precipitation of minerals from water solutions. We call such rocks chemical sedimentary rocks. They typically have a crystalline texture, partly formed during their original precipitation and partly when, at a later time, new crystals grow at the expense of old ones through a process called recrystallization. In some examples, crystals are coarse. In others, they are too small to see. Geologists distinguish among many types of chemical sedimentary rocks, primarily on the basis of composition.

Evaporites: the products of salt-water evaporation.  

In 1965, two daredevil drivers in jet-powered cars battled to be the first to set the land speed record of 600 mph. On November 7, Art Arfons, in the Green Monster, peaked at 576.127  mph; but eight days later Craig Breedlove, driving the Spirit of America, reached 600.601 mph. Travelling at such speeds, a driver must maintain an absolutely straight line; any turn will catapult the vehicle out of control. Thus, high-speed trials take place on extremely long and flat racecourses. Not many places can provide such conditions the Bonneville Salt Flats of Utah do. The salt flats formed when an ancient salt lake evaporated. Under the heat of the Sun, the water turned to vapour and drifted up into the atmosphere, but the salt that had been dissolved in the water stayed behind. Salt precipitation occurs where salt-water becomes supersaturated, meaning that it has exceeded its capacity to contain more dissolved ions. In supersaturated salt-water, ions bond to form solid grains that either settle out of the water or grow on the floor of the water body. Supersaturated salt-water develops where evaporation removes water from a water body faster than the rate at which new water enters. This process takes place in desert lakes and along the margins of restricted seas (figure above). For thick deposits of salt to form, large volumes of water must evaporate. Because salt deposits form as a consequence of evaporation, geologists refer to them as evaporites. The specific type of salt minerals comprising an evaporite depends on the amount of evaporation. When 80% of the water evaporates, gypsum forms; and when 90% of the water evaporates, halite precipitates. 

Travertine (chemical limestone).  

Travertine is a rock composed of crystalline calcium carbonate (CaCO3) formed by chemical precipitation from groundwater that has seeped out at the ground surface either in hot- or cold-water springs, or on the walls of caves. What causes this precipitation? It happens, in part, when the groundwater degasses, meaning that some of the carbon dioxide that had been dissolved in the groundwater bubbles out of solution, for removal of carbon dioxide encourages the precipitation of carbonate. Precipitation also occurs when water evaporates, thereby increasing the concentration of carbonate. Various kinds of microbes live in the environments in which travertine accumulates, so biologic activity may also contribute to the precipitation process. Travertine produced at springs forms terraces and mounds that are meters or even hundreds of meters thick, such as those at Mammoth Hot Springs (a in figure above). Travertine also grows on the walls of caves where groundwater seeps out (b in figure above). In cave settings, travertine builds up beautiful and complex growth forms called speleothems.

Dolostone.  

Another carbonate rock, dolostone, differs from limestone in that it contains the mineral dolomite (CaMg[CO3]2), which contains equal amounts of calcium and magnesium. Where does the magnesium come from? Most dolostone forms by a chemical reaction between solid calcite and magnesium-bearing groundwater. Much of the dolostone you may find in an outcrop actually originated as limestone but later changed into dolostone as dolomite crystals replaced calcite. This change may take place beneath lagoons along a shore soon after the limestone formed, or a long time later, after the limestone has been buried deeply.

Chert (replacement).  

A tribe of Native Americans, the Onondaga, once lived off the land in eastern New York State. Here, outcrops of limestone contain layers or nodules (lenses or lumps) of a black chert (a in figure above). Because of the way it breaks, the tribe’s artisans could fashion sharpedged tools (arrowheads and scrapers) from this chert, so the Onondaga collected it for their own toolmaking industry and for use in trade with other people. Unlike the deep sea (biochemical) chert described earlier, the chert collected by the Onondaga formed when cryptocrystalline quartz gradually replaced calcite crystals within a body of limestone long after the limestone was deposited; geologists call such material “replacement chert.” 

Chert comes in many colours (black, white, red, brown, green, gray), depending on the impurities it contains. Petrified wood is chert that forms when silica-rich sediment, such as ash from a volcanic eruption, buries a forest. The silica dissolves in groundwater, and then later precipitates as cryptocrystalline quartz within wood, gradually replacing the wood’s cellulose. The chert deposit retains the shape of the wood and the growth rings within it. Some chert, known as agate, precipitates in concentric rings inside hollows in a rock and ends up with a striped appearance, caused by variations in the content of impurities incorporated in the chert (b in figure above).